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Oxygen Isotope Signatures of the Apatite-Iron Oxide Ore at Grängesberg

Franz Weis

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Table of Contents

Abstract ...... 2 1. Introduction ...... 3 2. The local Geology of Grängesberg and the apatite-iron oxide ore ...... 4 2.1 The Bergslagen district ...... 4 2.2 Local Geology ...... 5 3. Oxygen isotopes ...... 6 3.1 Oxygen Isotopes and the δ-value ...... 6 3.2 Oxygen Isotope Fractionation ...... 6 3.3 Oxygen isotopes as geothermometer ...... 8 3.4 Oxygen isotopes as tracers ...... 8 4. Formation of the Ore Deposits at Grängesberg ...... 9 4.1 Origin by Magmatic Processes ...... 9 4.1.1 Direct magmatic deposition ...... 9 4.1.2 Ore Magma from Liquid Immiscibility ...... 9 4.2 Hydrothermal ore forming processes ...... 11 4.2.1 Magmatic waters ...... 11 4.2.2 Hydrothermal fluids from meteoric and sea water ...... 12 4.3 Other possible iron ore forming processes ...... 13 4.3.1 Hypothesis about the origin of apatite-iron oxide mineralization ...... 13 5. Expectations and assumptions ...... 14 5.1 Magmatic scenario ...... 14 5.2 Hydrothermal scenarios ...... 14 6. Oxygen isotope signatures of the Grängesberg deposit ...... 14 6.1 The samples from Grängesberg ...... 15 6.1.1 Sample preparation ...... 15 6.2 Oxygen isotope data ...... 15 6.3 Analytical methods and calculations ...... 18 6.3.1 Calculating δ18O-values for a possible ore forming magma ...... 18 6.3.2 Geothermometer calculations ...... 18 6.4 Results ...... 18 2

6.4.1 Results for the Geothermometer Calculation ...... 19 6.4.2 Results for ore magma modelling ...... 19 7. Discussion ...... 21 7.1 δ18O-Signatures across the Ore Body ...... 21 7.2.1 Magma fractionation ...... 21 7.2.2 Geothermometry calculations ...... 22 7.3 Comparing models and δ18O-results ...... 22 7.4 Comparison to δ18O-values of other apatite iron oxide deposits ...... 24 7.5 A possible formation of a mixed ore body ...... 25 8. Conclusion ...... 27 Acknowledgements ...... 28 References ...... 29

Abstract

The origin of apatite iron oxide ores, like the deposit at Grängesberg in the Bergslagen mining district, has been a subject of much discussion through the years. Some support a formation by hydrothermal fluids while others suggest that the ore is orthomagmatic, i.e. formed directly from a magma as suggested for the iron ore deposits of El Laco in Chile or the deposits in Kiruna, although also these two are still subjected to controversies. In 2009 sampling was done on drillcores through the Grängesberg ore. On these samples an oxygen isotope study on magnetite, quartz and whole rock samples from both the ore and its host rocks was conducted in order to obtain new knowledge about the ore forming processes. The data allowed modeling to simulate a possible origin by different magmas or hydrothermal fluids as well as a possible temperature of formation. In addition, the data set was compared to published oxygen isotope analyses of the possible magmatic iron ores of Kiruna and El Laco. The results of the Grängesberg analysis revealed that the ore in the area seem to have an origin from both magmatic and hydrothermal sources. 3

Oxygen isotope signatures of the apatite-iron oxide ore at Grängesberg

1. Introduction

Grängesberg is located in south central Sweden in the Bergslagen mining district. The iron oxide deposit at Grängesberg (Fig. 1) played a important role in Sweden’s mining history and economy since the 1800s, as the mine was the second largest iron ore producer in the country (Grängesberg Iron 2009) with a total production of 150 Mt between 1500 and 1989 (Allen et al. 2008). In addition, it is one of the few unique places with apatite-iron oxide ore, and makes up 40 % of the known iron content in Bergslagen (Allen et al. 2008). The apatite- iron oxide ore in the area has an outstanding Fe content of 40-63 % (Allen et al. 2008). However, falling iron ore prices caused the closure of the mine in 1989 (Grängesberg Iron 2009).

“Only iron can save us…” Fig. 1 The abandoned mine of Grängesberg. In the background the open pit infilled with water can be seen (www.bulkforum.net). This is what it says in the

poem “The cross of iron” by the German poet and writer Maximilian von Schenkendorf. What was important to the people of that time in the war against Napoleon, is equally important for us today. Iron is still essential for our economy and daily life. A modern world without iron is unthinkable and the demand for this highly valued element is still growing. Mining of iron began 1300 BC (de Beer et al. 1998) and has developed since. In 1945 global production was 100 Mt per year, increasing to 770 Mt in 1990 (de Beer et al. 1998) and in 2010 2400 Mt were produced (USGS 2011). Already in the 1990s the iron ore industry made up 15 % of the total industrial energy consumption worldwide and was thus the largest single sector with such a big demand (de Beer et al. 1998). Among the 15 biggest producers worldwide, where China is 4 the biggest with 900 Mt/year, Sweden is on place number 11 with 25 Mt/year. Due to this large demand and the promising possible output of 2.5 Mt/year (Grängesberg Iron 2009) and a reserve of 150 Mt (Allen et al. 2008) there are plans to reopen the Grängesberg mine. But besides finding new iron ore deposits or reopening closed mines in order to secure the needs, geologists are also interested in metallogenesis in order to gain information for future prospecting.

There are two main hypotheses of the formation of the apatite-iron oxide ore: by direct magmatic or hydrothermal processes (Corona-Esquivel et al. 2010; Frietsch and Perdahl 1994). These two processes are the most likely possible mechanisms when it comes to the formation of apatite iron ore (Juhani Ojala et al. 2008). Much research has been done on the origin of several iron ore deposits around the globe in order to unravel the ore forming processes. One of the methods used is the application of stable isotopes such as those of oxygen. The δ18O value in a rock or ore mineral can also be used as a signature and a geothermometer for the fluid or magma that was involved in the formation (e.g. Nyström 2008).

In this study oxygen isotope analysis has been conducted on the apatite-iron oxide ore at Grängesberg and the results are compared to oxygen isotope data from other similar deposits elsewhere.

2. The local Geology of Grängesberg and the apatite-iron oxide ore

2.1 The Bergslagen district

Grängesberg is located in the north-western part of the Bergslagen district. The mineralized rocks in the Bergslagen area consist mainly of the metamorphosed, Palaeoproterozoic Svecofennian volcano-sedimentary succession (Ripa et al. 2003). The volcanic to subvolcanic rocks are predominated of rhyolitic to dacitic composition and were mainly deposited in a submarine environment (Ripa 2001). However, also subordinate, intermediate and mafic volcanic rocks along with some chemical, epiclastic and organosedimentary rocks including carbonates occur at different stratigraphic levels in the volcanic pile (Ripa 2001). The basement of the Svecofennian rocks is not exposed but the presence of 2.7-1.95 Ga detrital- zircon in quartzites led to the assumption that the region is underlain by a grantitic basement (Allen et al. 1996). The base of the metavolcanic succession is exposed on Utö 5 island where a metasedimentary succession grades into the metavolcanic rocks (Allen et al. 1996). Large volumes of intrusive rocks were emplaced coeval with the volcanic activity (Ripa et al. 2003; Wilson et al. 1984). The rocks were all formed, deformed and metamorphosed during the Palaeoproterozoic Svekokarelian orogeny between 1.9 and 1.8 Ga (Stephens et al. 2009). In addition, the western part was also affected by the Sveconorwegian orogeny between 1.0-0.9 Ga (Stephens et al. 2009). The deformation of the metavolcaninc and metasedimentary rocks is predominantly expressed as steep, tight to isoclinal, doubly plunging F1 synclines (Allen et al. 1996).

The geological setting which led to the formation of the rocks in the Bergslagen district has been interpreted to be an extensional, back-arc basin inboard an active continental margin region (Allen et al.1996). This interpretation is supported by the occurrence of quartz-rich turbidites on Utö, quartz rich basement derived sandstones within the volcanic succession, the abundance of rhyolite and the overall chemistry of the volcanic rocks (Allen et al. 1996).

2.2 Local Geology

The western Bergslagen district is dominated by an 8 km thick metavolcanic succession (Allen et al. 1996). The host rocks of the apatite- iron oxide ore at Grängesberg are intermediate dacitic, andesitic as well as rhyolitic metavolcanic rocks which formed about 1.91-1.89 Ga (Stephens et al. 2009; Ripa et al. 2003). To the east and west these rocks were intruded by granitoids (Fig. 2) with mixing Fig. 2 The local Geology of Grängesberg. The red ellipse shows the position of the ore body (Map from SGU). and mingling features and younger, post-metamorphic granites and dolerite dykes (Ripa et al. 2003). The apatite-iron oxide deposit consists of 6 magnetite and hematite and occurs in general as a steeply south-eastward dipping (70°-80°) (Johansson 1910), lens-shaped body (Stephens et al.2009). The iron content ranges from 40- 63 % (Allan et al. 2008). The ore body can be followed for more than 900 m at the surface and it has a range in width between 50 m and 100 m (Johansson 1910).

3. Oxygen isotopes

3.1 Oxygen Isotopes and the δ-value

Both radiogenic and stable isotopes are a very useful tool in geosciences. They are used for age determination, estimating certain surface processes such as weathering and also for tracing climate changes and for petrogenesis. Significant for the latter are the two stable isotopes of oxygen, 16O and 18O. The heavier 18O isotope is also used as a tracer in order to identify the source of a certain rock or mineralization as well as a geothermometer (Rollinson 1993). The relative abundance of 18O is often presented as the δ18O-value given in parts per mil (‰). The δ18O-value is given by:

F3.1 δ18O = x 1000

(Rollinson 1993, Mason et al. 1982).

The standard value here is the Standard Mean Ocean Water value (short SMOW) which has a 18O/16O = 2005.20 +/- 0.43 (Hoefs 1997) derived from a hypothetical water sample. There are other standard values like the V-SMOW (Vienna-Standard Mean Ocean Water) or the PDB (Pee Dee Belemnite) but the one most commonly used is the SMOW-value (Hoefs 1997). The delta value shows whether the analyzed sample is either enriched (positive value) or depleted (negative value) in the heavier 18O relative to the standard (Rollinson 1993). The degree of enrichment or depletion is a result of isotope fractionation processes.

3.2 Oxygen Isotope Fractionation

Isotope fractionation is defined by Hoefs (Hoefs 1997) as “the partitioning of isotopes between two substances or phases of the same substance with different isotope ratios”. An example would be the oxygen isotope exchange between quartz and magnetite: 7

16 18 18 16 2Si O2 + Fe3 O4 = 2Si O2 + Fe3 O4 (Rollinson 1993).

This is also valid for fractionation between a rock and a fluid. By this exchange one substance can become enriched in the heavy isotope while the other remains depleted. The general rule is that the preferential site for the lighter isotope is with the weaker bond (Rollinson 1993). For example the heavier isotope is preferentially bonded to elements with higher ionic potential as they form stronger bonds. Thus quartz with the small, highly charged Si4+ ion will be enriched in 18O compared to magnetite with the larger, lower charged Fe2+ ion (Rollinson 1993). Fractionation between two species is shown in the form of the fractionation factor α which is given as

18 16 F3.2 α1-2 = where R is the O/ O in each species (Rollinson

1993, Mason et al.1982, Hoefs 1997, Javoy 1977). Furthermore it has been determined by experiment that

6 2 F3.2.1 1000 ln α1-2 = A(10 /T ) + B (Rollinson 1993, Hoefs 1997). A and B are constants called the thermometric coefficients which have been determined experimentally for particular mineral pairs (Javoy 1977) and T is the absolute temperature in Kelvin. B is often found to be zero and some values for A are shown in Table 3.2.

Table 3.2 Experimentally determined values for the thermometric coefficient A (Data after Chiba et al. 1989).

Thus the equation above shows that the oxygen isotope fractionation is temperature dependent. As the temperature increases the right side trends towards zero which means that α trends towards one. This infers that the fractionation is less at higher temperatures. Materials formed at lower temperature show larger variations in the δ18O values than do high temperature igneous rocks (Mason et al. 1982). Beside temperature the fractionation is marginally affected by the crystal structures and the content of ions of a species but is not 8 affected by pressure (Rollinson 1993). The impact of pressure on oxygen isotope fractionation is less than 0.1 % for pressures less than 20 kb.

3.3 Oxygen isotopes as geothermometer

Using equation F3.2.1, oxygen isotopes can be used as a geothermometer if the 18O/16O- ratios of the two species are known. However, it can also be linked to the δ-value, which is more commonly used as follows:

1. δ1 - δ2 ≈ 1000ln α1-2 for δ < 10 ‰

2. α1-2 = for δ > 10 ‰ (Rollinson 1993).

By inserting these equations into equation F3.2.1, the δ18O values of two minerals in a rock can be used to calculate the temperature at which the fractionation took place. This provides for example information about the temperature conditions under which a rock or mineral formed (Mason et al 1982).

3.4 Oxygen isotopes as tracers

A great variability of δ18O-values is found in nature because of the effect of isotope fractionation. The values vary about 100 ‰ in total. Fig. 3 shows the range for several δ18O- values for various rocks and water types. Oxygen Fig. 3 Variations of δ18O in natural waters and different rock types (Rollinson isotopes are a powerful 1993). tool for analyzing processes especially in the crust and upper mantle (Mason et al 1982). Since the δ18O signature, when the fractionation approaches equilibrium, is different for different settings, the oxygen isotopes can be used as a tracer for the origin of a specific mineral, rock or fluid including ore forming fluids (Hoefs 1997, Mason et al. 1982) as well as interactions and 9 mixing of different fluids. An example: Igneous rocks are usually enriched in the heavy isotope and show a δ18O between 6 and 10. Meteoric waters on the other hand, are heavily depleted and usually show δ18O-values of -5 to -10. If an igneous rock interacts with a hydrothermal fluid derived from meteoric water, fractionation leads to depletion in the heavy isotope by as much as 10 - 15 ‰ in the igneous rock (Taylor 1971). Thus, it is possible to detect the influence of a hydrothermal fluid on a magmatic rock.

4. Formation of the Ore Deposits at Grängesberg

There are two main hypotheses for the formation of the apatite-iron oxide deposits: direct magmatic and hydrothermal, which also could be applied to the mineralization in Grängesberg. In Part 4 these two possibilities will be discussed and it will be shown how an ore deposit could originate from either of these processes. The Information presented in the following paragraphs is mainly based on Guilbert and Park (1986) if not stated otherwise.

4.1 Origin by Magmatic Processes

4.1.1 Direct magmatic deposition

Formation of iron oxide mineralizations by magmatic processes can occur by direct deposition of an ore magma. An ore magma is any magma or magmatic fraction that can solidify directly from a melt as a mineralization. In most cases the magma must be ultra- basic i.e. enriched in mafic components. An iron oxide deposit could form in the simple way of fractional crystallization and crystal settling caused by cooling. The first mineral crystals that will separate out will be the ferromagnesian including magnetite along with chromite. Their crystallization temperature range lies from 1500° down to 800° C. When the magnetite crystals settle they may form a cumulate which then serves as an ore body (Park and MacDiarmid 1964). However, orthomagmatic oxide deposits can also form due to magma or liquid immiscibility.

4.1.2 Ore Magma from Liquid Immiscibility

Liquid immiscibility (Fig. 4) is another possible mechanism for ore formation (oxide magma formation) and probably more common than direct deposition from a parent magma. This term stands for the process when a given magma due to cooling or crystallization can no longer exist in a stable state and thus splits into two melts with different compositions 10

(Roedder 1978). The composition of the newly generated melts depends on the composition of the parent magma (Roedder 1978). If one considers cooling of a basaltic magma that starts to crystallize the first formed silicates, that are less silicic than the parental magma, the composition of the evolved melt Fig. 4 The process of liquid immiscibility. As a basaltic magma will shift to be more felsic. Due to undergoes crystallization and differentiation, the silica rich melt may separate into two liquids depending on the enrichment of this the amount of sulfur and either oxygen or sulfur. The point of separation is shown by the red dot (Diagram after Guilbert and Park 1986). oxygen increases in the residual melt, which will be considered as a silicate melt enriched in those elements. Due to cooling or further enrichment the melt may no longer be capable of keeping the elements in solution. At this point liquid immiscibility occurs which leads to the generation of a separate oxygen or sulfur enriched melt. Normally only one type of melt will be generated. Two melts may form at a rather late stage of the magma evolution, but also in a very early stage if the magma is enriched in oxygen or sulfur from the start. In this case two liquids are present in the magma chamber and among them partition of certain elements occurs. In the case of separate oxygen melt siderophile/oxyphile elements such as Cr, V, P, Ti and most important Fe would partition into sulfur/oxide melt while other lithophile elements stay behind in the silicate liquid. In such a way an iron and oxygen enriched melt would form, from which an iron oxide mineralization could be generated.

The generated ore magmas can either erupt at the surface or solidify at depth to form an intrusion. If a partly crystallized magma containing the ore fluids is exposed to differential external stresses the ore magma may be separated from the crystal mush in the process called filter pressing, in which a mat of crystals compacts under its own weight and expels interstitial melt (Sisson and Bacon 1999). The ore bearing melt can thus intrude into the surrounding country rock and form a magmatic injection deposit (Fig. 5). Ore bodies formed from direct magmatic processes are rather common (Park and MacDiarmid 1964). 11

Fig. 5 Ore formation by magmatic processes (Diagram after Guilbert and Park 1986).

4.2 Hydrothermal ore forming processes Fig.5 Ore formation by magmatic processes (Diagram after Guilbert). The formation of a mineralization by hydrothermal fluids involves the circulation of hot water or more specified hot water solutions from which the ore minerals can precipitate (Marshak 2008). Hydrothermal fluids can be of various origins. They may be the result of magmatic processes, be meteoric water percolating at depth or of metamorphic origin (Marshak 2008, Park et al 1964, Guilbert et al 1986). Metamorphic waters are however rarely assumed to be responsible for ore formation and are thus not discussed further.

4.2.1 Magmatic waters

Hydrothermal fluids which are generated by magma are called juvenile waters or magmatic

waters. Magma always contains volatiles including H2O. Granitic to andesitic magmas contain between 1 and 7 wt. % water (Clemens 1984), some even up to 15 wt. %. The water content increases with differentiation of the magma, i.e. a more evolved magma has higher water content. This is related to the ongoing differentiation and crystallization in the magma chamber in which the lighter residual liquid and the volatiles accumulate at the top of the chamber. These magmatic waters are essential in the transportation of metals like the

transition metals, alkali earths, halogens and other volatiles such as CO2. The metals are

carried in solution in the form of metal halogen complexes such as for example FeCl2. With 12 ongoing differentiation the water content in the residual liquid increases and eventually reaches a point at which the water separates from the magma. This was shown experimentally and presented by Smith in 1948. However, this process requires certain temperature and pressure conditions. If there is a rupture in the surrounding country rock the magmatic water can be lost in a similar way as CO2 escapes from a bottle of sparkling water when the cap is removed. The water pressure can also increase to a level exceeding the lithostatic pressure causing an explosion-breccitaion. Another possible formation of magmatic water bubbles is by exsolution. In this case the water cannot escape the magma chamber. The escaping, metal-bearing water can intrude the country rock and form mineralized fault veins or breccia columns. If the juvenile water reaches a sea-floor it will result in a black smoker or as a thermal spring in a subaerial environment. The formation and separation of magmatic waters is associated with temperatures of the range 500° - 300° C.

4.2.2 Hydrothermal fluids from meteoric and sea water

Metal-bearing waters, mineralizers, can also be generated by the heating of meteoric water or sea water. Water percolating downwards will be heated up due to the geothermal gradient or by an igneous intrusion nearby (Chernicoff et al. 2002). For this, the country rock must have a certain permeability. As the meteoric water assumes thermal equilibrium with its surrounding it also moves towards chemical equilibrium. This means that as the temperature of the water increases the amount of dissolved minerals from the surrounding country rock will gradually increase (Park and MacDiarmid 1964; Guilbert and Park 1986). A rock or even a magma in contact with the hydrothermal fluid are leached of ore-forming elements like iron (Keays 1987, Lagerblad et al. 1985). In such a way a natural brine is formed which may be enriched in iron. Iron rich rocks or magmas, i.e. mafic or ultra-mafic rocks, will be more viable to enrich the fluid. From such a natural brine iron oxide can precipitate if the fluid cools or there are pressure changes or if the acidity in the fluid drops (Marshak 2008). Ore minerals precipitated from these types of hydrothermal fluids are often found in cracks as vein deposits or along faults and bedding planes (Chernicoff 2002, Lianxing et al. 1993). Repeated leaching and precipitation of already formed iron oxide mineralizations will result in further enrichment and the generation of high grade ores (Ünlü et al. 1995, Lianxing et al. 1993). 13

4.3 Other possible iron ore forming processes

Iron oxide mineralization can also be generated by sedimentary processes and occur as placer deposits (Marshak 2008) or formed by metamorphic processes but these are not considered for the apatite-iron oxide deposits. The previously described processes involve much more detail than shown in Fig. 6 but it provides a brief overview of ore formation by magmatic and hydrothermal fluids.

4.3.1 Hypothesis about the origin of apatite-iron oxide mineralization

Examples of iron oxide deposits that are interpreted to have originated from direct magmatic processes are the magnetite-hematite deposit of El Laco in Chile (Park et al. 1964), the apatite-iron oxide ore body of Kiruna in Sweden (Frietsch et al. 1994; Nyström 2008; Park et al. 1964) and the La Perla iron oxide ash fall deposit in Mexico (Corona-Esquivel et al. 2010). But even these deposits are still under discussion like the one in Grängesberg. There is the hydrothermal school proposed by Paràk (1975, 1984), Hitzman et al. (1992), Sheets (1997), Rhodes and Oreskes (1994, 1999), Sillitoe and Burrows (2002) and there is the magmatic school supported by Nystöm et al (2008), Frietsch (1978) and Geijer (1931, 1967). However, uncontroversed hydrothermal iron oxide deposits are for example the Hamersley iron ore in Western Australia (Barley et al. 1999) and the ore bodies in the Yangtze Valley

District in China (Lianxing et al. 1993).

Fig. 6 Summary of ore forming processes by magmatic and hydrothermal processes (Diagram after Guilbert and Park 1986).

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5. Expectations and assumptions

A few assumptions and expectations can be drawn for the Grängesberg apatite-iron oxide ore related to whether the origin is magmatic or hydrothermal. These two scenarios are discussed and assumptions are made for what the oxygen data of the mineralization are expected to show in each case.

5.1 Magmatic scenario

Magmas of andesitic and intermediate composition have a δ18O-value between 5.7 and 10 ‰ (Fig. 3). A mineralization formed by direct magmatic processes in these rocks would have a similar δ18O-signature. The same would also account for juvenile waters, which fall into the same range. An example for this would be the Olympic Dam deposit in Australia. The δ18O- value of the iron oxides was found to be ~ 10 ‰ and thus the source from which it formed was interpreted to be magmatic (Hunt 2005). However, the fractionation between the ore minerals and the magma has to be taken into account and the δ18O-value of the ore mineral will be lower than the magma. According to Sheppard (1969) a magnetite formed from a hot magmatic aqueous fluid will have a δ18O-value about 5 ‰ lower than the fluid itself and Taylor (1967) states that most igneous magnetites fall into a range of 1-4 ‰ when the equilibrium with the magma is reached.

5.2 Hydrothermal scenarios

The δ18O-values for high-T hydrothermal fluids are below 5.7 ‰. This would be the case for a hydrothermal fluid originating both from sea water and meteoric waters. Thus, ore deposits that were derived from these fluids are expected to also have a hydrothermal signature and should not lie in the magmatic range. Since the fractionation is temperature dependent, the δ18O of a hydrothermal ore deposit should lie below the magmatic range for high-T and above the magmatic range for low-T hydrothermal processes.

6. Oxygen isotope signatures of the Grängesberg deposit

Oxygen isotope analysis was carried out on both whole rock samples and mineral phases from both the ore body and the host rocks of the apatite-iron oxide deposit at Grängesberg. The analysis was carried out at Cape Town University, South Africa using a DeltaXP dual inlet gas source mass spectrometer. 15

6.1 The samples from Grängesberg

The samples collected for the oxygen isotopes were selected from two drillcores (drillcore 690 and 717) that transect the ore body at the Export field. The samples include the ore and meta-volcanic host rocks of andesitic, dacitic and rhyodacitic compositions. From these oxygen isotope analyses for whole rock samples, quartz, magnetite and apatite were carried out.

6.1.1 Sample preparation

The two drillcores were logged and sampled in 2009 by the Swedish Geological Survey (SGU). The selected samples were prepared by Barbara Kern at Uppsala University. The rock samples were crushed into small pieces using a hammer. After crushing the aggregated samples were separated into different grain size fractions which were washed in distilled water for 2 minutes and then dried in an oven. Picking of the different mineral grains was done under a binocular microscope. About 400 mg of each mineral species were picked and the crystals were chosen by their size and purity. The whole rock samples were crushed and milled.

6.2 Oxygen isotope data

In total from the two drillcores 38 δ18O-values were obtained from whole rock, magnetite, apatite and quartz. Of these analyses 28 are from drillcore 717 and 15 from 690. Tables 6.1 and 6.2 give a description of the samples and the obtained δ18O-values. Due to the mineral composition of the rocks a full set of data was not possible to reach for all samples.

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Sample Description Whole Rock Quartz Magnetite Apatite KES090048 Foliated, locally phyllosilicate rich volcanic rock 6.91 8.01 KES090052 Gneissic granite with magnetite crystals 9.08 3.47 KES090054 Foliated, veined, Bt-rich volcanic, partly sköl-like 7.05

KES090056 Grey,porphyritic volcanic rock with several mm large feldspar phenocrysts 6.93

KES090059 Feldspatic-porphyritic,light coloured, felsic to intermediate volcanic-subvolcanic rock 6.63 KES090061 Bt-Kl-banded volcanic to subvolcanic, partially altered rock 6.2 KES090065 Mt veinlets/bands in grey, banded volcanic rock 4.89 5.80 1.29 KES090071 Massive mt-ore section with almost no apatite or silicates 1.92 KES090073 Apatite-rich section, Mt-ore 1.34 5.50

KES090076 Massive He-ore, with some apatite and magnetite 3.4

KES090077 Mt-blastic He ore, apatite-bearing 0.18 KES090080 He-ore with magnetite blasts 1.56 1.08 KES090081 Alteration zone, phyllosilicate rich, contact to He-blastic Mt-ore 1.53 KES090083B Mt-banded volcanic rock 2.78 KES090086 Mt-banded grey, felsic to intermediate volcanic rock. Close to ore contact 1.69 KES090087 Fine grained, foliated, felsic-intermediate volcanic to subvolcanic rock with mt bands 7.09 5.95

KES090088 Densely, finely banded Mt-skarn-volcanic rock 5.64 3.44

KES090089 Foliated, fine grained, felsic-intermediate volcanic rock 7 KES090069 Massive Mt-ore, with spots of silicate minerals, no visible apatite 1.41 3.04

Table 6.1 The sample descriptions and δ18O-values of the 717 drillcore.

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Sample Description Whole Rock Quartz Magnetite KES090003 Slightly fsp porhyritic dark volcanic/subvolcanic 6.20 7.40 KES090004 Alteration zone, amph, bt, fsp, locally py-bearing 6.97 KES090007 Bt-veined, dark volcanic rock Locally sköl-like 5.20 KES090009 Apatite-rich banded ore 6.00 2.24 KES090014A Reddish porphyritic volcanic to subvolcanic rock 8.95 KES090039 Phyllosilikate rich alterd volcanic rock 3.82 3.69 KES090040 Phyllosilicate-veined volcanic rock, schlieren-like 5.47 KES090044 Mt-blastic intermediate to basic volcanic rock 6.10 -0.42 KES090045 Bt rich, schistose, intermediate to basic volcanic rock with some Mt blasts. 6.47 KES09008A Mt-ore, immediately near contact to host-rock 1.93 KES090037 Mt ore section, close to host rock contact. 1.38

Table 6.2 The sample descriptions and δ18O-values of the 690 drillcore. 18

6.3 Analytical methods and calculations Table 6.2 The sample description and δ18O of the 690 drillcore. The ore forming fluid leaves an oxygen isotope imprint on the ore minerals. However, it has also been noticed that fractionation occurs between the solid and liquid phase. In order to take this fractionation into account and to use the data for modeling a few different Table 6.2 The dataset for the 690 drillcore. calculations were carried out using the equations in chapter 3.

6.3.1 Calculating δ18O-values for a possible ore forming magma

The magnetite δ18O-values can be used to reconstruct a model for potential ore forming magma. This can be done if the fractionation (1000lnα) between a magma type and a mineral is known. If the δ18O-value of a specific mineral and the fractionation are known, a theoretical δ18O-value for the potential ore forming magma can be obtained. In order to do

so the known values have to be inserted in δ1 - δ2 ≈ 1000ln α1-2 and the equation has to be rearranged to solve the unknown. The fractionation values for magnetite for rhyolitic and andesitic magma are -4.67 and -3.97 respectively (Zhao and Zheng 2002). The calculation for a rhyolitic magma is:

18 18 -4.67 ‰ = δ Omagnetite - δ Otheoretical magma

In which the theoretical magma value is the unknown. With this equation a range of potential δ18O-value for rhyolitic and andesitic magmas are obtained which can then be analyzed and compared to the expected range of these magma types and the whole rock values of the samples.

6.3.2 Geothermometer calculations

The equation F3.2.1 gives the temperature at the time of the fractionation. The calculation is done for the samples where both magnetite and quartz data exists. The value for the geothermometer coefficient A is taken from Chiba (1989).

6.4 Results

For the first set of calculations all the magnetite values were used in order to obtain models for a possible ore magma. Overall, 16 possible values were obtained and are presented in tables 6.4 and 6.5. A potential range of magma δ18O for rhyolites and andesites was determined (Table 6.3) by the rule that the δ18O increases by 0.2 ‰ with every 5 wt. % 19

increase of the SiO2 content with respect to the mantle values. This gave the following theoretical ranges:

Magma SiO2 Content Theoretical δ18O in ‰ Crustal Contamination (+2 ‰) Andesite 52 % -63 % 6-6.22 6-8.22 Dacite 63 %-69 % 6.22-6.46 6.22-8.46 Rhyolite > 69 % >6.46 > 6.46

Table 6.3 Theoretical source magma δ-values

However, these values are only based on fractional crystallization and some crustal contamination had to be assumed (+ 2 ‰) which would extend the range.

6.4.1 Results for the Geothermometer Calculation

The geothermometer calculation is done for two samples where the mineral pair “quartz- magnetite” exists. The calculations and results are shown below.

For the granite sample: KES090052

18 18 δ O-value Quartz: 9.08 δ O-value Magnetite: 3.47 δ1 - δ2 : 5.61

 5.61 = 6.29(106/T2)

 T=

 T= 1058 K = 785°C

For the metavolcnic sample: KES090065

18 18 δ O-value Quartz: 5.80 δ O-value Magnetite: 1.29 δ1 - δ2 : 4.51

 4,51 = 6.29(106/T2)

 T=

 T= 1180 K = 907°C

6.4.2 Results for ore magma modelling

The two tables below show the results of the calculations for andesitic and rhyolitic magmas as well as a comment regarding a comparison with the expected magma range and the whole rock value for the samples. 20

Sample δ18O Magma Value from Andesite fractionation Comments on value Value Whole Rock KH09005 Ba mgt 0.33 4.30 too low KES090052 mgt 3.47 7.44 fits KES090081 mgt 1.53 5.50 too low KES090083B mgt 2.78 6.75 fits KES090088 mgt 3.44 7.41 fits 5.64 KES090069 mgt 3.04 7.01 fits 1.41 KES090071 mgt 1.92 5.89 too low KES090086 mgt 1.69 5.66 too low KES09008A mgt 1.93 5.90 too low KES090009 mgt 2.24 6.21 fits 6.00 KES090037 mgt 1.38 5.35 too low KES090039 mgt 3.69 7.66 fits 3.82 KES090044 mgt -0.42 3.55 too low 6.10 KES090065 mgt 1.29 5.26 too low KES090077 mgt 0.18 4.15 too low KES090080 mgt 1.08 5.05 too low 1.56

Table 6.4 The ore magma model for the andesite magma fractionation.

Sample δ18O Magma value from Rhyolite fractionation Comments on value Value Whole Rock KH09005 Ba mgt 0.33 5.00 too low KES090052 mgt 3.47 8.14 fits KES090081 mgt 1.53 6.20 too low KES090083B mgt 2.78 7.45 fits KES090088 mgt 3.44 8.11 fits 5.64 KES090069 mgt 3.04 7.71 fits 1.41 KES090071 mgt 1.92 6.59 fits KES090086 mgt 1.69 6.36 too low KES09008A mgt 1.93 6.60 fits KES090009 mgt 2.24 6.91 fits 6.00 KES090037 mgt 1.38 6.05 too low KES090039 mgt 3.69 8.36 fits 3.82 KES090044 mgt -0.42 4.25 too low 6.10 KES090065 mgt 1.29 5.96 too low KES090077 mgt 0.18 4.85 too low KES090080 mgt 1.08 5.75 too low 1.56

Table 6.5 The ore magma model for the rhyolite fractionation.

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7. Discussion

7.1 δ18O-Signatures across the Ore Body

The δ18O-values from the ore deposit are expected to have an imprint from the ore formation. The data show that in general all the magnetite values and most of their whole rock samples have a δ18O-value below 5.7 ‰. The first impression of this result is that the ore is of hydrothermal origin. However, since fractionation Fig. 7 δ18O vs. position of the samples in the drillcore 690. has to be taken into account this is questioned and tested by doing the modeling. Figures 7 and 8 show the δ18O result of the drillcores against distance in meters.

7.2 Results from modeling

7.2.1 Magma fractionation

The calculations for magma fractionation modeling give Fig. 8 δ18O vs. position of the samples in the drillcore 717. some interesting results. These show that with the theoretical ore forming magma’s δ18O-values would fall into several ranges using the fractionation values from Zhao and Zheng (2002). For the andesite fractionation six out of sixteen magnetite samples show that they might have an imprint of their δ18O-value from the fractionation with an andesitic magma. For the rhyolite magma fractionation model eight samples fall into the magmatic range. For both, the samples that 22

did not fall into the magma fractionation model are considered to have a hydrothermal origin (Fig. 9).

Fig. 9 The source magma modeling result.

7.2.2 Geothermometry calculations

The results from geothermometry calculations show expected magmatic temperatures for the granite with T= 785°C and T= 907°C for the metavolcanic rock with magnetite veins. In fact these temperatures would fit to a rhyolitic/granitic magma (temperature range 900° - 700° C).

7.3 Comparing models and δ18O-results

The results from the calculations and modeling are promising but they have to be compared with the analytical data, particularly the whole rock samples. The theoretical magma values have to be compared with the obtained analytical results and the calculated temperatures have to be compared with the model. The dataset from the modeling includes six samples that include δ18O-values from both magnetite and whole rock. Three of these samples fall into the magmatic range (Tables 7.4 and 7.5).

Magma Value from Andesite Comments Value Whole Sample Sample Description δ18O fractionation on value Rock KES090088 mgt Magnetite Ore 3.44 7.41 fits 5.64 KES090009 mgt Apatite rich banded Ore 2.24 6.21 fits 6 KES090039 mgt Metavolcanite 3.69 7.66 fits 3.82 KES090044 mgt Metavolcanite -0.42 3.55 too low 6.1 KES090069 mgt Magnetite Ore 3.04 7.01 fits 1.41 KES090080 mgt Magnetite Ore 1.08 5.05 too low 1.56

Table 7.4 The comparison of δ18O-values from whole rock to theoretical magma source for andesite fractionation 23

Magma value from Rhyolite Comments Value Whole Sample Sample Description δ18O fractionation on value Rock KES090088 mgt Magnetite Ore 3.44 8.11 fits 5.64 KES090009 mgt Apatite rich banded Ore 2.24 6.91 fits 6 KES090039 mgt Metavolcanite 3.69 8.36 fits 3.82 KES090044 mgt Metavolcanite -0.42 4.25 too low 6.1 KES090069 mgt Magnetite Ore 3.04 7.71 fits 1.41 KES090080 mgt Magnetite Ore 1.08 5.75 too low 1.56

Table 7.5 The comparison of δ18O-values from whole rock to theoretical magma source for rhyolite fractionation

Sample KES090088 (marked yellow in table 7.4) is from a laminated metavolcanic rock within the ore and has magnetite that shows a theoretical magmatic origin. The whole rock value also falls within the mantle value (5.7 +/- 0.3 ‰) which possible for andesitic or rhyolitic magmas (Fig. 3). This leads to the assumption that the ore was generated by an ore magma. If a slight hydrothermal alteration/overprint, which would lower the δ18O-value of the host rock, is assumed, a perfect fit for the magmatic origin is gained. Sample KES090069 is from a massive magnetite part of the ore and the modeling fits for a magmatic origin. The whole rock value is however low. But again a hydrothermal overprint could be assumed, which may affect the δ18O-value drastically.

On the other hand, sample KES090080 which is a magnetite ore with hematite blasts (marked blue in table 7.4) seems to have originated from hydrothermal fluids. The magma modeling clearly indicates this by showing a theoretical magma δ18O-value of 5.05 which is too low for being magmatic. Also the whole rock δ18O-value of this sample is very low (1.56 ‰) indicating a hydrothermal origin or overprint as suggested by the hematite content.

The three remaining samples are metavolcanic rocks and an ore sample (KES090009) and the modeling for the latter, an apatite-rich banded ore, shows that the magnetite is supposed to have formed from a magma also indicated by the whole rock value of 6 ‰. In KES090044, an intermediate volcanic rock with magnetite blasts, the whole rock value shows a magmatic value (6.1 ‰) but the magnetite is supposed to be of hydrothermal origin (-0.42 ‰). The magnetite probably formed as a secondary mineral during hydrothermal alteration associated with the volcanic activity. The whole rock value is just at the lower boundary of the assumed andesite δ18O-range, so hydrothermal alteration can be assumed. KES090039, a 24 phyllosilicate altered volcanic rock, with a whole rock value of 3.82 ‰ again contains magnetite that probably crystallized from a magma according to the model, and was thus not affected by the later hydrothermal event altering the host rock.

Only for those six samples, described above and presented in Table 7.5, whole rock data is available to be compared with the modeling. The origin of the other samples can only be assumed by the calculations. Another possible magmatic ore sample is KES090083B, a magnetite banded volcanic rock, with a magnetite δ18O-value of 2.78 ‰ and KES090071, an apatite-banded magnetite ore, with 1.91 ‰ for the magnetite which however would only work if it had its origin by a rhyolitic magma. The same holds for sample KES09008A, a magnetite ore. Sample KES090077, a magnetite blastic hematite ore, and KES090037, a magnetite ore, both in turn represent two more hydrothermally affected parts of the ore body by theory with a δ18O-value of 0.18 ‰ and 1.38 ‰ respectively. The remaining samples included in the modeling are magnetite bearing metavolcanic rocks. Again, the results indicate that some may have formed directly from a magma or may be the result of later stage hydrothermal alteration.

The two samples for which the fractionation temperature was calculated were KES090052, a granite, and KES090065, a metavolcanic rock with magnetite veins. Both indicated a fractionation occurring at magmatic temperatures. The oxygen isotope data from the granite and the modeling showed that this magnetite has a magmatic origin. This fits with the temperature of 785°C. The second sample, KES090065, is a metavolcanite and the fractionation between the quartz and the magnetite occurred at around 907°C. This is a magmatic temperature. On the other hand, the δ18O of the magnetite together with the modeling indicate that this magnetite in the rock did not form from magma. A possibility could be that the magnetite actually formed by a hydrothermal process, maybe a mineralizer from the magma or by later hydrothermal overprint or the two minerals simply had not reached equilibrium.

7.4 Comparison to δ18O-values of other apatite iron oxide deposits

The δ18O-result from the apatite-iron oxide deposit at Grängesberg can be compared with the results from other similar ore deposits for which an oxygen isotope anaylsis is available. Nyström et al. (2008) deals with the oxygen isotope analysis of the El Laco deposit in Chile which they compare to the ore in Kiruna. Three diagrams show oxygen isotope values for 25 magnetite from Kiruna, El Laco and the Chilean Iron Belt. By comparing magmatic textured magnetite in Nyström et al. (2008) to the values from Grängesberg it can be shown that the values fall into the same range of δ18O-values. A special similarity can be observed with the values from the El Laco deposit where the δ18O-value of magnetite falls into a range between 2 – 4 ‰ which supports the interpretation and the modeling that the apatite-iron oxide ore at Grängesberg, which falls into the same range, is also magmatic. However, there is still the hydrothermal school for the El Laco deposit (Rhodes and Oreskes 1999).

Fig. 10 shows several magnetite δ18O-values from several iron ore deposits whose formation is under discussion. In part A of the diagram the similarities with the El Laco deposit clearly stand out as well with the Chilean Iron Belt deposits. The lower boundary for magmatic derived magnetites on the diagram was taken to be the magnetite that would have formed and be in equilibrium with a MORB (5.7 ‰). The upper boundary was taken to be a magnetite in equilibrium with a rhyolitic magma with a δ18O-value of 10 ‰. Data plotting in between these values is considered to be of an orthomogmatic origin, meaning that it crystallized purely from a magma (Fig. 10). Data plotting to the left of the diagram below the lower limit of 5.7 ‰ is seen to have been overprinted by or originated from hydrothermal processes including juvenile and/or magmatic waters. To determine the exact hydrothermal processes that were involved in the magnetite formation is beyond the aim of this study. In total a magmatic origin would account for six magnetite samples out of which four are from the main ore. The remaining ten magnetite samples are of hydrothermal origin of which seven are ore bodies. In order to obtain a bigger picture of which type, hydrothermal or magmatic, is in majority more samples would have to be analyzed.

7.5 A possible formation of a mixed ore body

The formation of such an ore body with both magmatic and hydrothermal signatures can be possible in a volcanic system in a subduction zone. A possible scenario is a volcano which was fed by a magma enriched in oxides (may it be due to liquid immiscibility or fractionation). This would lead to either an eruption of oxide lava forming a magmatic ore body or to the mineralization of apatite-iron oxide ore at depth. Parts of the ore may have originated from magmatic waters. Other hydrothermal processes are initiated by meteoric or sea water percolating downwards into the country rock being heated by the magma. This causes leaching of the country rocks and maybe some ore bodies and leads to mineralization 26 of apatite-iron oxide ore from a hydrothermal fluid. Both types of ore forming processes are indicated in the Grängesberg samples.

Fig. 10 Magnetite δ18O-values from different iron ore deposits. Part A serves as a comparison of the Grängesberg magnetites with magnetites from other deposits. Part B shows the magnetites assumed to be of purley magmatic origin.

The present orientation of the ore body as a steeply dipping sheet is due to later tectonic events (Fig. 11) and unrelated to the ore forming processes. In order to reconstruct the exact 27 formation of the ore body at Grängesberg more research would have to be done and would involve more than just and oxygen isotope analysis.

Fig. 11 A possible formation of a mixed apatite-iron oxide ore body. The present orientation is due to tectonic events.

8. Conclusion

The oxygen isotope analysis of the apatite-iron oxide ore at Grängesberg provided insight about the origin of the ore. Over the years it has been a controversy whether the formation was due to hydrothermal or purely magmatic processes (Johansson 1910, Magnusson 1938). Evaluation of the δ18O-values of the Grängesberg samples provides that in fact both processes have been involved. Modeling of a possible ore forming magma showed that direct magmatic origin can be assigned to six out of the sixteen samples as they fall into the magmatic window. This was also supported by the comparison of the magnetite δ18O-values to those from other magmatic ore deposits where the values were in the same range. The other ten Grängesberg samples fall into the hydrothermal field. The question what types of hydrothermal processes were involved in the formation is out of the scope of this study and would require more investigations. Also more sampling and oxygen isotope analysis would be necessary to reveal which ore forming process formed the main ore body. An ore body of a mixed hydrothermal and magmatic type would be possible in a subduction zone at an active continental margin setting like the Bergslagen region in the Paleoproterozoic. Oxygen isotope data in this study indicate a formation of the apatite-iron oxide ore at Grängesberg not solemnly by hydrothermal or magmatic processes, but a mixture of both.

28

Acknowledgements

I want to thank my supervisors Dr. Karin Högdahl and Prof. Valentin R. Troll at Uppsala University, who gave me the opportunity to work on this project, for all their help and support. Also I want to thank Prof. Erik Jonsson from the Swedish Geological Survey (SGU; Uppsala University), who did the drillcore logging and who provided the samples for the project, Dr. Chris Harris from the University of Cape Town for carrying out the oxygen isotope analysis and providing the data and Barbara Kern at the Technical University of Luleå for the sample preparation.

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